Optical properties of chromophoric (CDOM) and fluorescent dissolved organic
matter (FDOM) were characterized in the Nordic Seas including the West
Spitsbergen Shelf during June–July 2013, 2014, and 2015. The CDOM absorption
coefficient at 350 nm, aCDOM(350) showed significant interannual
variation (T test, p< 0.00001). In 2013, the highest average
aCDOM(350) values
(aCDOM(350) = 0.30 ± 0.12 m−1) were observed due
to the influence of cold and low-salinity water from the Sørkapp Current (SC) in
the southern part of the West Spitsbergen Shelf. In 2014, aCDOM(350)
values were significantly lower (T test, p< 0.00001) than in 2013
(average aCDOM(350) = 0.14 ± 0.06 m−1), which was
associated with the dominance of warm and saline Atlantic Water (AW) in the
region, while in 2015 intermediate CDOM absorption (average
aCDOM(350) = 0.19 ± 0.05 m−1) was observed.
In situ measurements of three FDOM components revealed that
fluorescence intensity of protein-like FDOM dominated in the surface layer of
the
Nordic Seas. Concentrations of marine and terrestrial humic-like DOM were
very low and distribution of those components was generally vertically
homogenous in the upper ocean (0–100 m). Fluorescence of terrestrial and
marine humic-like DOM decreased in surface waters (0–15 m) near the
sea ice edge due to dilution of oceanic waters by sea ice meltwater. The
vertical distribution of protein-like FDOM was characterized by a prominent
subsurface maximum that matched the subsurface chlorophyll a maximum and
was observed across the study area. The highest protein-like FDOM
fluorescence was observed in the Norwegian Sea in the core of warm AW. There
was a significant relationship between the protein-like fluorescence and
chlorophyll a fluorescence (R2= 0.65, p< 0.0001,
n= 24 490), which suggests that phytoplankton was the primary source
of protein-like DOM in the Nordic Seas and West Spitsbergen Shelf waters.
Observed variability in selected spectral indices (spectral slope
coefficient, S300–600, carbon-specific CDOM absorption
coefficient at 254 and 350 nm, SUVA254, aCDOM*(350)) and
the nonlinear relationship between CDOM absorption and the spectral slope
coefficient also indicate a dominant marine (autochthonous) source of CDOM
and FDOM in the study area. Further, our data suggest that
aCDOM(350) cannot be used to predict dissolved organic carbon
(DOC) concentrations in the study region; however the slope coefficient
(S300–600) shows some promise in being used.

The rapid reduction of summer sea ice in the Arctic Ocean in the past decades
has various repercussions on the structure and functioning of the Arctic
marine system, forcing changes in physics, biogeochemistry, and ecology of
this complex oceanic system (Meier et al., 2014). One of the most significant
consequences of observed rapid Arctic Ocean transition is an increase in the
primary productivity of the Arctic Ocean (Arrigo et al., 2008), which could
potentially contribute to increased production of autochthonous (marine)
dissolved organic matter (DOM) in ice-free and under-ice waters. The sea ice
is also a source of autochthonous DOM and its chromophoric (colored) subfraction, CDOM (e.g., Granskog et al., 2015a;
Anderson and Amon, 2015; Retelletti-Brogi et al., 2018). However, dissolved organic carbon
(DOC)
produced by ice algae has a limited effect on overall organic carbon mass
balance in the Arctic Ocean, as melting of 1 m of sea ice would
negligibly change DOC concentration in the top 50 m of the water column, assuming an
averaged DOC content in the ice of 100 µMol C (Anderson and Amon,
2015). Simultaneously, response of terrestrial ecosystems to temperature
increase will accelerate permafrost thaw and increase the riverine discharge,
resulting in more allochthonous (terrestrial) DOM being released into the
Arctic Ocean (Amon, 2004; Stedmon et al., 2011; Anderson and Amon, 2015;
Prowse et al., 2015, and references therein). Terrestrial DOM plays a
considerable role in the carbon budget of the Arctic Ocean (Findlay et al.,
2015; Stein and Macdonald, 2004), especially in coastal waters and the
continental shelf with a large inflow of terrestrial DOM, which constitutes
80 % of total organic carbon delivered by Arctic rivers (Stedmon et al.,
2011).

The optically active DOM fraction called CDOM represents light-absorbing molecules (Coble, 2007;
Nelson and Siegel, 2013; Stedmon and Nelson, 2015). Once entered or produced
in surface waters of the Arctic Ocean, CDOM has a significant influence on
heating of the uppermost ocean layer and its stratification (Pegau, 2002;
Hill, 2008; Granskog et al., 2007, 2015b). Particularly in the absence of sea
ice, light absorbed by CDOM in the visible part of the spectrum limits the light
available for photosynthetic organisms (Arrigo and Brown, 1996) but also
shields marine ecosystems from potentially harmful ultraviolet radiation by
strongly absorbing electromagnetic radiation in the UVB and UVA bands (Erickson
III et al., 2015). CDOM is also an important substrate in photochemical
reactions contributing to direct remineralization of organic carbon,
production of bioavailable low-molecular-weight DOM but also formation of
reactive oxygen species that could potentially be toxic to marine organisms
(Mopper and Kieber, 2002; Kieber et al., 2003; Zepp, 2003). The
mineralization by photochemical reactions or microbes of DOM, both terrestrial
and marine, is a crucial but still insufficiently quantified mechanism in the
Arctic carbon cycle (e.g., Osburn et al., 2009). Despite the importance of
CDOM, studies on its distribution, properties, and transformation in the
Arctic Ocean and its marginal seas are still limited, partly by their
remoteness and seasonal accessibility.

The North Atlantic sector of the Arctic Ocean is a region with a complex
interaction between inflowing warm and highly productive AW entering
the Arctic and cold and fresh Polar Surface Water (PSW) exiting the Arctic Ocean.
Recent studies have reported intensification of AW inflow
into the Arctic Ocean (Walczowski, 2014; Polyakov et al., 2017; Walczowski et
al., 2017), further highlighting the importance of the European sector of the
Arctic Ocean to better understand the complex interactions between inflowing
AW and PSW. Optically these waters are contrasting, especially with
respect to CDOM (Granskog et al., 2012; Pavlov et al., 2015; Stedmon et al.,
2015) and FDOM (Jørgensen et al., 2014; Gonçalves-Araujo et al.,
2016). In the absence of sea ice, favorable vertical mixing conditions and
sufficient levels of solar radiation make it a very productive and important
region from an ecosystem and socioeconomic standpoint, thus ensuring
motivation for ongoing studies of the complex marine system in the area
(Skogen et al., 2007; Olsen et al., 2009; Dalpadado et al., 2014). In the context
of ongoing and further anticipated intensification of Atlantic Ocean inflow
to the Arctic Ocean, a description of processes and factors controlling
CDOM and FDOM properties and distribution could be used to better predict future
changes associated with CDOM in the areas upstream of the AW
inflow region, to estimate glacial meltwater (Stedmon et al., 2015), and
to trace water masses (Gonçalves-Araujo et al., 2016).

A number of occasional synoptic surveys of CDOM and optical properties have
been conducted in the different regions of the European Arctic Ocean and
concentrated on the western part of the Fram Strait influenced by polar water
outflow with the East Greenland Current (EGC) (Granskog et al., 2012; Pavlov et al., 2015;
Gonçalves-Araujo et al., 2016). The CDOM distribution in the area
influenced by AW was reported by Stedmon and Markager (2001) in the central
part of the Greenland Sea, and by Granskog et al. (2012) and Pavlov et
al. (2015), who presented CDOM and particulate absorption distribution along
a
transect across the Fram Strait at 79∘ N. Hancke et al. (2014)
studied the
seasonal distribution of the CDOM absorption coefficient
(aCDOM(λ)) in an area across the Polar Front in the
central part of the Barents Sea. Seasonal studies on CDOM contribution to
overall variability in inherent optical properties (IOPs) reported on
sea ice (Kowalczuk et al., 2017) and in the water column during a spring
under-ice phytoplankton bloom north of Svalbard (Pavlov et al., 2017). In
this study we aimed to present variability in CDOM and FDOM optical
properties in a large area spanning parts of the Barents, Norwegian, and
Greenland seas (particularly focusing on the West Spitsbergen Shelf) over
a period of 3 consecutive years (2013–2015) and understand the role of
(i) large-scale ocean circulation patterns and water mass distribution and
(ii) phytoplankton productivity as controlling factors on CDOM and FDOM
distribution.

2.1 Study area

Observations were carried out in the framework of the long-term observational program
AREX, conducted since 1987 by the Institute of Oceanology, Polish Academy of
Sciences, Sopot, Poland, and covered the area of water mass exchange between
the North Atlantic Ocean and the Arctic Ocean (Fig. 1). The Norwegian,
Barents, and Greenland seas, called the Nordic Seas, represent a crucial
component of the Northern Hemisphere climate system due to two contrasting
water masses and their contribution to the heat and salt exchanges between
the North Atlantic and the Arctic Ocean (Walczowski, 2014; Schlichtholz and
Houssais, 1999a, b). The warm and salty AW are carried
northward by the North Atlantic Current (NAC), which further splits into two
major branches. The Norwegian Current (NC) flows into the Barents Sea as the
Barents Sea branch, while the West Spitsbergen Current (WSC) heads north
along the eastern flank of the Fram Strait. The EGC
flows south along the western side of the Fram Strait and carries cold and low-salinity PSW and sea ice (Fig. 1) (e.g., Schlichtholz and
Houssais, 2002). The East Spitsbergen Current (ESC) could also affect the region
with transformed polar water originating from the northeastern Barents Sea
(Sternal et al., 2014). The main ESC branch flows southward along the coast of
Spitsbergen and its extension is the SC, which influences the
West Spitsbergen Shelf. The remaining part of polar water from the Barents Sea
flows southwestward along the eastern slope of the Spitsbergenbanken
towards Bear Island as the Bjørnøya Current (Loeng, 1991) in the
Norwegian Sea and the Barents Sea border. Presence and extensiveness of polar
water from the Barents Sea depends on favorable wind conditions affecting the
magnitude and the exchange with the AW inflow (Nilsen et al., 2015;
Walczowski, 2014).

Optical measurements and water sampling were conducted during three summer
Arctic expeditions (AREX) onboard R/V Oceania in 2013, 2014, and 2015
(AREX2013, AREX2014, and AREX2015, respectively) (Table 1). In situ FDOM
fluorescence measurements were conducted in 2014 and 2015. AREX expeditions
covered the Norwegian Sea with a main section along the border between the
Norwegian Sea and the Barents Sea (sampled in late June to early July 2014
and 2015). The area of the western and northern Spitsbergen shelf was
investigated in July of 2013–2015 (Fig. 1), along sections spanning from
shelf towards the sea ice edge. The westernmost and northernmost sampling
stations north of 76∘ N, shown in Fig. 1, correspond to the sea ice
edge position in July in the given year.

Table 1Dates of AREX expeditions, and number of samples or number of in
situ vertical profiles of CDOM, DOC, chlorophyll a (Chla) inherent
optical properties (IOPs), chlorophyll a fluorescence
(IFChla), and FDOM fluorescence.

2.2 Sample collection and processing

Water samples for determination of CDOM absorption, chlorophyll a, and DOC
were collected with a Sea-Bird SBE32 Carousel Water Sampler equipped with Niskin
bottles, an SBE 911plus conductivity–temperature–depth (CTD) probe (SBE 9plus CTD unit and SBE
11plus Deck Unit), and WET Labs ECO chlorophyll fluorometer. Samples were
collected at three depths: near the surface, ca. 2 m depth, at the
chlorophyll a maximum, which was usually located between 15 and 25 m depth,
and below the chlorophyll a maximum, between 50 and 70 m. The exact position
of chlorophyll a maximum depth was estimated from the vertical profile of
chlorophyll a fluorescence during the CTD downcast. During AREX2013 water
samples for CDOM absorption measurements were immediately filtered in two
steps: first through acid-washed GF/F filters, and second through
acid-washed Sartorius 0.2 µm pore size cellulose membrane filters
to remove finer particles. In 2014 and 2015 CDOM samples were filtered
directly from rosette Niskin bottles through a Millipore Opticap XL4 Durapore
filter cartridge with nominal pore size 0.2 µm into acid-washed
200 mL amber glass bottles. The cartridge filter was kept in 10 % HCl
solution and was rinsed with ultrapure Milli-Q and sample water before
collecting CDOM samples. In 2013 and 2015 collected unpreserved water samples
for determination of CDOM absorption were stored onboard R/V
Oceania in the dark, at a temperature of 4 ∘C, and were
transferred after the cruise to a land-based laboratory for spectroscopic
measurements. In 2014, all spectroscopic measurements for the determination of
CDOM absorption were carried out in the laboratory onboard R/V Oceania,
immediately after collection. Samples for determination of DOC concentration
were collected the same way as CDOM samples. Water that passed through
0.2 µm filters was collected into pre-cleaned 40 mL glass vials
(certified pre-cleaned sample vials, Sigma-Aldrich) and acidified with a drop
of concentrated 38 % HCl. Acidified samples were stored onboard the ship
in the
dark, at a temperature of 4 ∘C, and were transferred after the cruise
to a land-based laboratory for measurements.

Water samples for the determination of chlorophyll a concentration were
filtered immediately after collection under low vacuum on Whatman (GE
Healthcare, Little Chalfont, UK) 25 mm GF/F filters. Filter pads with
particulate material retained on them were immediately deep frozen in a
freezer and thereafter stored at −80 ∘C prior to analyses.

2.3 CDOM absorption

Before spectroscopic scans were conducted, the temperature of the CDOM
absorption samples was increased to room temperature. CDOM absorption for
AREX2013 and AREX2015 was measured using a double-beam PerkinElmer LAMBDA
650 spectrophotometer in the spectral range 240–700 nm, in the laboratory
at the Institute of Oceanology, Polish Academy of Sciences in Sopot, Poland.
Measurements of the CDOM absorption samples collected during AREX2014 were
carried out onboard the research vessel, using a double-beam PerkinElmer LAMBDA
35 spectrophotometer in the same spectral range as in 2013 and 2015. The
10 cm quartz cuvette was chosen for all measurements and the reference was
fresh ultrapure water. Absorbance A(λ) spectra were transformed to
the CDOM absorption coefficients, aCDOM(λ) (m−1),
according to

(1)aCDOMλ=2.303⋅A(λ)/L,

where 2.303 is the natural logarithm of 10, A(λ) is the corrected
spectrophotometer absorbance reading at a specific wavelength (λ), and
L is the path length of optical cell in meters (here 0.1 m).

The slope coefficient of the CDOM absorption spectrum, S, between 300 and
600 nm was derived using Eq. (2) and was implemented in MATLAB R2011b by
adopting a nonlinear least-squares fit with a trust-region algorithm
(Stedmon et al., 2000; Kowalczuk et al., 2006):

(2)aCDOM(λ)=aCDOM(λ0)e-S(λ0-λ)+K,

where λ0 is a reference wavelength (here 350 nm), and K is a
background constant representing any possible baseline shifts not due to CDOM
absorption. Simultaneous calculation of three parameters:
aCDOM(350), S, and K was performed according to Eq. (2) in the
spectral range between 300 and 600 nm by nonlinear regression. CDOM
absorption coefficient values are also included at two other wavelengths,
aCDOM(375) and aCDOM(443), to enable direct comparison
of our results with previously published studies. In 2014 the range of the
spectral slope coefficient had to be reduced to 300–500 nm due to spectra
disturbances over 500 nm in the data set from the western and northern
Spitsbergen shelf. To assess the effect of the narrower spectral range on
spectral slope coefficient calculations, we calculated slopes for both
spectral ranges in 2013 and 2015. On average, the spectral slope coefficient in
the spectral range 300–500 nm was higher by 1.76 µm−1
relative to S300–600. Calculated average bias was deduced from
S300–500 calculated in 2014 to comply with the 2013 and 2015 data
sets. A linear regression model was used on log-transformed CDOM absorption
spectra for spectral slope coefficient calculations at the spectral range
275–295 nm, S275–295.

2.4 Chlorophyll a concentration

Filters pads containing suspended particles (including pigments) were used
for determination of the chlorophyll a concentration for all AREX cruises.
Pigments were extracted at room temperature in 96 % ethanol for 24 h.
Spectrophotometric determination of chlorophyll a concentration,
Chla
[mg m−3], was performed with two spectrophotometers: UV4–100 (Unicam, Ltd)
and a PerkinElmer LAMBDA 650 in 2013 and 2014–2015, respectively. The
optical density (absorbance) of pigment extract in ethanol was measured at
665 nm. Background signal was corrected in the near-infrared region of the electromagnetic spectrum (750 nm):
ΔOD= OD(665 nm) − OD(750 nm). Subsequently,
conversion of absorbance to chlorophyll a was carried out according to the following
equation (Strickland and Parsons, 1972; Stramska et al., 2003):

(3)Chla=(103⋅ΔOD⋅VEtOH)/(83⋅Vw⋅l),

where 83 (dm3 (g cm)−1) is the chlorophyll a specific absorption
coefficient in 96 % ethanol, Vw (dm3) is the volume of
filtered water, VEtOH (dm3) is ethanol extract volume, and
l is the path length of the cuvette (here 2 cm).

2.5 DOC concentration

DOC measurements were performed with a “HyPer+TOC” analyzer (Thermo Electron
Corp., the Netherlands) using UV persulfate oxidation and nondispersive
infrared detection (Sharp, 2002). Potassium hydrogen phthalate was used as a
standard addition measurement method for each sample in triplicate.
Consensus reference material (CRM) supplied by Hansell Laboratory from the
University of Miami was analyzed as a quality control of DOC concentrations.
The methodology provided sufficient accuracy (average recovery 95 %;
n= 5; CRM = 44–46 µM C; our
results = 42–43 µM C) and precision represented by a relative
standard deviation (RSD) of 2 %.

The carbon-specific CDOM absorption coefficient at 350 nm, aCDOM∗(350) (m2 g−1), was determined as the ratio of the
CDOM absorption coefficient at a given wavelength aCDOM(350) to
the DOC concentration (Eq. 4):

(4)aCDOM∗(350)=aCDOM(350)DOC,

where DOC is expressed in milligrams per liter.

The carbon-specific UV absorption coefficient (SUVA) is defined as the UV
absorbance of the water sample at a specific wavelength normalized for DOC
concentration [mg L−1] (Weishaar et al., 2003). SUVA
(m2 gC−1) at 254 nm (SUVA254, Eq. 5) is an indicator of
aromaticity of aquatic humic substances and was calculated as

Vertical profiles of IOPs, FDOM, and
chlorophyll a fluorescence together with conductivity, temperature, and
pressure were measured at all stations from the surface down to 200 m depth
using an integrated instrument package consisting of an ac-9 plus
attenuation and absorption meter (WET Labs Inc., USA), a WETStar CDOM
fluorometer (WET Labs Inc., USA), a microFlu-chl chlorophyll a fluorometer
(Trios GmbH, Germany), and a Sea-Bird SBE 49 FastCAT
CTD probe (Sea-Bird Electronics, USA).

Spectral light absorption, a(λ), and beam attenuation, c(λ),
coefficients were measured at nine wavelengths (412, 440, 488, 510, 532, 555,
650, 676, and 715 nm). The ac–9 plus calibrations were performed
regularly. After cleaning with ultrapure water, stability instrument
readings were inspected with in-air measurements. The required correction of
absorption signal for scattering was performed with the so-called proportional
method by which zero absorption is estimated at 715 nm (Zaneveld et al., 1994).
Subtraction of absorption coefficients from attenuation coefficients
determined volume scattering coefficient, b(λ). The excitation channel
and maximum emission of light detector of the microFlu-chl
chlorophyll a fluorometer were set at 470 and at 686 nm, respectively.
Recorded chlorophyll a fluorescence intensity signals,
IFChla, were reported as analog voltage output in the range
0–5 V DC. The instrument setup is described in detail in Granskog et
al. (2015b).

FDOM was measured using a three-channel WET Labs WETStar fluorometer equipped
with two laser LEDs that excited the water sample inside the flow-through
quartz cell at 280 and 310 nm, and two detectors to measure emission
intensity at 350 and 450 nm. Such construction allowed for combinations of
three channels with distinct excitation–emission features in specific peak
areas as given in Coble (1996): Channel 1 (CH1), ex. = 310 nm and
em. = 450 nm, represents marine ultraviolet humic-like peak C and marine
humic-like peak M; Channel 2 (CH2), ex. = 280 nm and em. = 450 nm,
represents UVC terrestrial humic-like peak A; and Channel 3 (CH3),
ex. = 280 nm and em. = 350 nm, represents the protein-like
tryptophane peak T (Fig. S1 in the Supplement). ICHn is the
fluorescence intensity at a particular channel, where n denotes the channel
number from 1 to 3. Recorded ICHn could be transformed from raw
instrument counts into either the quinine sulfate equivalent (QSE) units, or
particular compound concentrations with factory calibration curves.
Application of the factory calibration curves, especially the blank ultrapure
water readings offset, resulted in negative values for ICH1 and
ICH2. Therefore, we reported fluorescence intensities acquired
from the WETStar fluorometer in raw counts (RC) corrected for a noticeable
but small drift. This offset was determined as the difference in any
ICHn, between initial measurements in July 2014 in the depth
range 100–150 m, at salinity > 34.9, and temperature
T> 0 ∘C and measurements repeated in the same salinity and
temperature range during the field campaign in 2015. The water salinity and
temperature characteristics at the chosen depth range were typical for the
core of AW inflow, which is characterized with stable values of spectral
absorption (measured with an ac–9 plus attenuation and absorption meter),
negligible chlorophyll a, and very low background CDOM absorption levels
(Sławomir Sagan, personal communication, 2017). Therefore, we assume that
any differences in raw WET Labs WETStar three-channel fluorometer readings
between measurements in 2014 and 2015 resulted from instrument drift, and the
offset between the years has been subtracted from florescence intensity
values at each channel measured in 2015.

2.7 Classification of water masses

Water masses were classified according to Rudels et al. (2005) based on
potential temperature (Θ), potential density (σθ), and
salinity (S). The original classification definitions are derived for Fram
Strait (Rudels et al., 1999) and categorization used in Rudels et al. (2002,
2005) considers mainly the EGC, the area of the Yermak Plateau and Storfjorden
located on the east coast of Spitsbergen. To adjust the classification to the
broader area of the Nordic Seas including the Atlantic part (Norwegian and Barents
seas), some modifications have been introduced (see Table S1 in the
Supplement).

The epipelagic layer of the Nordic Seas is dominated by AW and PSW and
waters formed in the mixing process and local modifications (precipitation,
sea ice melt, riverine runoff, and surface heating or cooling) of these
two water masses. AW masses were usually characterized by potential
temperature and density thresholds defined by Rudels et al. (2005)
(Table S1). To better distinguish AW from PSW, we added a third criterion:
any water mass classified as PSW (Rudels et al., 2005) with salinity higher
than S> 34.9 has been considered to be AW. The salinity criterion equal
to 34.9 is widely used in the literature (Swift and Aagaard, 1981;
Schlichtholz and Houssais, 2002; Walczowski, 2014) and eliminates the classification ambiguity of Rudels et
al. (2005) caused by modification of AW by local
sources of fresh water. Part of AW (except PSW warm, PSWw)
included waters with a density below σθ=27.7 kg m−3
(marked in Fig. 3 with dashed isopycnal line) used by Rudels et al. (2005) as
a threshold value between AW and PSW. Lower density of waters of Atlantic
domain with high salinity (> 34.9) is predominantly caused by high
temperatures and cannot be referred to as PSW, the lower density of which is
attributed to lower salinity. PSW is defined as Θ≤0∘C and σθ≤27.7 kg m−3. The
temperature of PSW is usually negative; however, positive temperatures
(3–5 ∘C) can be observed during summer (Swift and Aagaard, 1981).
Warmer PSWw has been considered here with the same σθ≤27.7 kg m−3 criterion and Θ> 0 ∘C (Rudels et
al., 2005), due to summer season measurements and higher temperatures of low-salinity surface waters in the eastern Fram Strait. Furthermore, PSWw was also
limited to the uppermost 50 m of the water column with S≤34.9. The
water mass with similar temperature–salinity (TS) characteristics to PSWw but slightly different
ranges was referred to in the literature for Arctic Surface Water, ASW
(e.g.,
Pavlov et al., 2015; Gonçalves-Araujo et al., 2016), but due to the
dominance in the area of water originating from the Atlantic Ocean the name PSWw
from Rudels et al. (2005) classification is used. We could find Arctic
Atlantic Water (AAW) in our data set as a result of the mixing process of AW and
PSW, in the range of 0 <Θ≤2∘C and
27.7 <σθ≤27.97 (Rudels et al., 2005). Arctic
Intermediate Water (AIW) was defined as Θ≤0.3∘C and
27.97 <σθ, σ0.5≤30.44 (Rudels et al., 2005)
and included measurements taken at the greatest depth in this study.

Figure 2Surface distribution of temperature, salinity, and
aCDOM(350) in 2013–2015 (a–c, respectively). Plots
were created with the use of Ocean Data View (Schlitzer, 2016).

3.1 Interannual and spatial variability in CDOM properties in surface waters with relation to hydrography

Spatial distribution of temperature, salinity, and aCDOM(350) in
surface waters of the West Spitsbergen Shelf and Norwegian Sea shows considerable
variation among years (Fig. 2). In 2013, the West Spitsbergen Shelf was
under the influence of cold and low-salinity waters from SC. The impact of this
current together with possible terrestrial runoff (the highest
aCDOM(350) values were observed at Spitsbergen fjord entrances)
was reflected in high aCDOM(350) (average
0.47 ± 0.26 m−1) for coastal waters on the West Spitsbergen
Shelf. Lower values of aCDOM(350) were observed in the PSWw
(average
0.33 ± 0.17 m−1) from coastal areas and in the warm and salty AW
from the WSC (average 0.28 ± 0.07 m−1). The lowest CDOM absorption
(average 0.25 ± 0.05 m−1) in 2013 was observed at the northernmost
and northeasternmost stations influenced by low-salinity PSW affected by sea ice
meltwater.

A quite different situation was observed in 2014 (Fig. 2b). The spatial
extent of AW was distinctly wider, as shown by temperature and salinity
distributions. The higher proportion of AW over the West Spitsbergen Shelf in
2014 was confirmed by the temperature and salinity time series in the top 200 m water layer (Walczowski et al., 2017). This large volume of AW influenced
CDOM absorption, which was lowered to half of the values (average
aCDOM(350) = 0.15 ± 0.06 m−1) compared to 2013.
In addition, mean aCDOM(350) values around 0.1 ± 0.03 m−1
were observed in the northern Spitsbergen shelf in the area affected by sea
ice melting (within the salinity range of 31.4–33.9).

In 2015, SC and ESC branches originating from the Barents Sea were
pronounced, as indicated by lower temperature and salinity, Fig. 2c,
resulting in elevated aCDOM(350) values on the West Spitsbergen
Shelf and along the section from Sørkapp down to 74∘ N and near
Bjørnøya Island. In 2015 AW was characterized by a low CDOM concentration
(aCDOM(350) average 0.17 ± 0.02 m−1) in contrast to PSW
observed north of Svalbard (average
aCDOM(350) = 0.27 ± 0.05 m−1).

Summary statistics of the variability in aCDOM(350),
aCDOM(443), S275–295, S300–600, aCDOM*(350), and SUVA254 in different
water masses in a given year are provided in Table 2. The highest
aCDOM(350) was observed in 2013 (Table 2) when CDOM absorption in
AW and PSW was similar (average
aCDOM(350) = 0.28 ± 0.07 m−1). CDOM absorption in
PSWw was higher and was characterized by the greatest variability (average
aCDOM(350) = 0.32 ± 0.16 m−1; min–max:
0.15–0.9 m−1 CV = 50 %; CV: coefficient of variation). In 2014 aCDOM(350)
values were almost 2 times lower compared to other summer seasons (Table 2).
In 2014 79 % of all samples were classified as AW (average
aCDOM(350) = 0.14 ± 0.05 m−1), which corresponded to the highest temperature, widespread AW
distribution, and lack of apparent influence by SC waters. Less than 15 %
of
samples represented PSWw (average
aCDOM(350) = 0.14 ± 0.05 m−1) (Table 2). In 2015
we observed intermediate aCDOM(350) values in AW and PSWw
(Table 2) with the highest values in PSW and AAW (PSW:
aCDOM(350) = 0.26 ± 0.09 m−1; AAW:
aCDOM(350) = 0.25 ± 0.06 m−1).

The spectral slope coefficient is often inversely nonlinearly related to the
CDOM absorption coefficient (Stedmon and Markager, 2001; Stedmon et al.,
2003; Kowalczuk et al., 2006; Meler et al., 2016). S275–295 and
S300–600 were lowest in 2013 and highest in 2014, with
intermediate values in 2015 (Table 2). The carbon-specific CDOM absorption
coefficient aCDOM*(350) was significantly lower
(p< 0.000001, T test) in 2014 compared to 2013 and 2015. The
values of SUVA254 were most diverse in 2013 whereas the greatest
variability in AW (min–max: 0.64–9.23 m2 gC−1) was observed in
2014. In 2014 and 2013 average values of SUVA254 for the whole season were
similar, around 1.7 m2 gC−1 (Table 3); however average values in
AW and PSWw were higher in 2013 and 2014, respectively (Table 2). In 2015
average SUVA254 values were similar within identified water masses and
low variation (±0.15 m2 gC−1) among different waters was
observed. The interannual variability in SUVA254 was insignificant
(p= 0.89, T test) between 2013 and 2014; however the average
SUVA254 values observed in 2015 were significantly different
(p< 0.002, T test) than in 2013 and 2014 (Table 2).

The average DOC concentration in the study area was highest in 2013
(80.69 µmol L−1) and
decreased significantly (p< 0.000001, T test) year by year (Table 3)
to 67.64 µmol L−1 in 2015. The average chlorophyll a concentration
was lowest in 2013 (0.87 mg m−3), almost doubled in 2014
(1.58 mg m−3), and decreased by 12 % in 2015 (1.39 mg m−3),
relative to the previous year.

3.2 Optical properties of different water masses

All measured salinity and temperature values are presented in the
TS diagram as a function of depth (Fig. 3a) to
visualize water masses sampled during the AREX2013, AREX2014, and AREX2015
campaigns. The majority of measurement represented characteristics of AW that
covered all depth ranges. The second water mass represented in our data set
was low-density PSWw (σθ≤27.7 kg m−3), which was
observed above 50 m depth. The smallest fraction of data points belonged to
PSW, which was aggregated in the subsurface 20–70 m depth range, and AAW, which
was encountered within the 50–100 m depth range (Fig. 3a). To visualize the
distribution of DOM properties within classified water masses we have chosen
the fluorescence intensity of the marine humic-like DOM (ICH1),
fluorescence intensity of the protein-like DOM (ICH3), and CDOM
absorption aCDOM(350). The highest ICH1 values were
observed in PSW and the lowest in PSWw (Fig. 3b). Humic-like FDOM in AW was
characterized by a large dynamic range and both low (320 RC) and high
values (> 360 RC) were observed (Fig. 3b). In the case of ICH3
the highest values were observed in PSW, PSWw mid depth (15–50 m, which can
be associated with chlorophyll a maximum), and in part of AW, which was
separated from PSWw (upper part: T> 0, σθ≤27.7,
S> 34.9). The lowest ICH3 values were observed in AW
(lower part: 27.7 <σθ≤27.97) and in PSWw, where
σθ≤26.5 (Fig. 3c). There was a large variability and no
consistent trends in distribution of aCDOM(350) values in
different water masses in the study area, as shown in the TS diagram
(Fig. 3d). The distribution of fluorescence intensity of the terrestrial
humic-like DOM (ICH2) and SUVA254 in the TS diagram is
shown in the Supplement (Fig. S2).

3.3 Vertical distribution of FDOM components

The instrumental in situ synchronous IOP measurements enabled us to resolve FDOM
distribution with better resolution, compared to coarser discrete water
sampling of CDOM. Representative vertical profiles of temperature, salinity,
FDOM, and chlorophyll a fluorescence are shown in Fig. 4. Differences in the
vertical distribution of salinity and temperature (Fig. 4a, b) were observed
at sampling stations located near the sea ice edge (black stars), where a
cold and fresher surface layer (typically 5–10 m deep; classified as PSWw)
was present. The salinity at stations located in the core of AW
(green circles) and at the southwestern Spitsbergen shelf (red circles) was
uniform in the upper 100 m (Fig. 4a, b). There was very little spatial and
vertical variation in humic-like FDOM (ICH1 and
ICH2). The only exception was the slightly higher, but still
vertically homogenous, distribution of humic-like FDOM observed at stations
near the Spitsbergen coast in 2014 (red dots; Fig. 4c, d).

The vertical distribution of protein-like FDOM (ICH3, Fig. 4e)
was very similar to distribution of chlorophyll a fluorescence
(IFChla, Fig. 4f) and the total non-water absorption coefficient
at 676 nm (atot−w(676), Fig. 4g). All three parameters had a
strong subsurface maximum at the depth range between 10 and 30–40 m and
similar spatial distribution. The surface values for these three parameters
were higher than values below the maximum (40 m) for profiles in the AW
(green and red symbols). Near the ice edge, however, stations were
characterized by lower values in the surface layer, comparable to the values
below 40 m, likely due to dilution of FDOM and Chla by sea ice meltwater at the very surface. The atot−w(676) vertical profiles in
AW were different, with elevated values throughout the whole upper layer
(0–30 m depth), which dropped sharply to a background level below the
subsurface chlorophyll a maximum.

3.4 Relationship between chlorophyll a and protein-like FDOM

The qualitative correspondence between fluorescence of protein-like FDOM and
chlorophyll a fluorescence intensity (Fig. 4) has been quantitatively
confirmed by regression analysis. A significant positive relationship between
ICH3 and IFChla was found in both 2014 and 2015
(R2= 0.65, p< 0.0001, n= 24 490; Fig. 5a). The
relationship was stronger in 2014 (R2= 0.75, p< 0.0001,
n= 17 700; blue line in Fig. 5a), when broader influence of AW
was observed (Walczowski et al., 2017), than in 2015 (R2= 0.45,
p< 0.0001, n= 7290; red line in Fig. 5a).

The same relationship was confirmed using data from discrete water samples. A
statistically significant relationship between ICH3 and
Chla
values was found in both years, and the determination coefficient for the
combined data set was R2= 0.36 (p< 0.0001) (Fig. 5b). There
was higher correlation observed between ICH3 and Chla values
in 2015 compared to 2014 (Fig. 5b). Higher dispersion between FDOM
fluorescence intensity measured in situ and chlorophyll a measured in water
samples could be a result of the time lag between instrumental measurements
and water collection that reached up to 1.5 h. The IOP instruments'
deployment was usually performed simultaneously with CTD downcast, while water
sample collection was performed during CTD rosette upcast, which was
significantly delayed especially at deep water stations (at sampling stations
located at a water depth > 1000 m). Observed higher protein-like FDOM
values per chlorophyll a concentration unit could be explained by
phytoplankton physiological response due to higher water temperature observed
in 2014 and consequently more efficient extracellular DOM release. This
physiological effect is evident in relationships between chlorophyll a
fluorescence and atot−w(676). In 2014 phytoplankton were more
fluorescent at the same absorption level (Fig. S3).

Figure 5Relationship between chlorophyll a fluorescence
(IFChla) and fluorescence of the protein-like component
(ICH3)(a) and relationship between fluorescence of the
protein-like component (ICH3) and chlorophyll a concentration
from discrete water samples (b) in the upper 200 m of the water column
in 2014 and 2015. Set of linear regression functions, correlation coefficient
(R), coefficient of determination (R2), p value, and number of
samples (n) are presented in Fig. 5.

4.1 Variability in and spectral properties of CDOM in the Nordic Seas

The highest CDOM absorption in the Arctic has been observed in coastal
margins along the Siberian Shelf in the Laptev Sea, close to the Lena River
delta (aCDOM(440) = 2.97 m−1 salinity close to 0)
(Gonçalves-Araujo et al., 2015) and in Laptev Sea shelf water at the
surface (aCDOM(443) > 1 m−1, salinity < 28)
(Gonçalves-Araujo et al., 2018) and at the coast of the Chukchi Sea and
southern Beaufort Sea influenced by riverine inputs of the Yukon and
Mackenzie rivers (aCDOM(440) > 1 m−1,
salinity < 28) (Matsuoka et al., 2011, 2012; Bélanger et al., 2013).
Exceptionally high CDOM absorption has also been observed in the southern
part of Hudson Bay near river outlets with
aCDOM(355) > 15 m−1, at a salinity close to 0 (Granskog
et al., 2007). Pavlov et al. (2016) reported aCDOM(350) of up to
10 m−1 at a salinity of 21 in surface waters of the White Sea.
Terrestrial CDOM from the Siberian Shelf has been diluted and
aCDOM(440) decreased to ca. 0.12 m−1 at salinities of 32.6
(Gonçalves-Araujo et al., 2015) and transported further toward the Fram
Strait by the Transpolar Drift, being gradually diluted or removed (Stedmon
et al., 2011; Granskog et al., 2012). In the Transpolar Drift and the central
Arctic Ocean, CDOM absorption in surface waters was dominated by terrestrial
sources with observed aCDOM(443) values varying between ∼ 0.15 m−1, at salinities close to ±27 (Lund-Hansen et al., 2015),
and ∼ 0.5 m−1 at a salinity range from 26.5 to 29.5
(Gonçalves-Araujo et al., 2018). Dilution also effectively decreased CDOM
absorption in the western Arctic Ocean, and average CDOM absorption in the
Chukchi Sea and Beaufort Sea was
aCDOM(440) = 0.046 m−1, at salinities > 32.3
(Matsuoka et al., 2011, 2012; Bélanger et al., 2013). The influence of
transformed AW generated in the Barents and Norwegian seas had impacted
aCDOM(443) values in the Beaufort Gyre and Amundsen and Nansen
basins, causing its decrease below 0.2 m−1 as reported by
Gonçalves-Araujo et al. (2018).

The reported lower range of aCDOM(350) observed in AW during
AREX2014 (2014: 0.14 ± 0.06 m−1) is in good agreement with data
from the eastern part of Fram Strait at the 79∘ N section reported by
Granskog et al. (2012), Stedmon et al. (2015), and Pavlov et al. (2015) and
with data reported by Hancke et al. (2014) south of the Polar Front in the
Barents Sea. Kowalczuk et al. (2017) observed similar aCDOM(350)
values
north of Svalbard. Higher values of CDOM absorption in AW observed in 2015
were within the published variability range (Pavlov et al., 2015; Hancke et al.,
2014; Kowalczuk et al., 2017). The highest aCDOM(350) values in AW in
2013, 0.28 ± 0.07 m−1 (Table 2), were similar to Hancke et al.
(2014) north of the Polar Front in the Barents Sea. Very low values of
aCDOM(443) aligned with previous reports: in the core AW in the Greenland Sea measured during TARA expedition in 2013 (Matsuoka et
al., 2017), in the eastern Fram Strait (Pavlov et al., 2015), in the
Barents Sea (Hancke et al., 2014), and north of Svalbard (Kowalczuk et al.,
2017). It should be underlined that data comparison could be biased by the number
of observations, as this study documented aCDOM(350) and
aCDOM(443) statistics based on a significantly higher number of
samples and wider spatial coverage compared to the sources cited above.

The AW inflow with the WSC is an extension of the NAC originating from the
Atlantic Ocean, and CDOM absorption presented in this study was comparable
with values found in the North Atlantic Ocean (Kowalczuk et al., 2013;
Kitidis et al., 2006). In contrast, values of absorption coefficients were
2 times higher in Norwegian coastal waters, which are influenced by the Lofoten
Gyre, and presumably by terrestrial runoff as reported by Nima et al. (2016).

Despite lower-salinity and lower-temperature, CDOM optical properties in PSW
in this study did not differ significantly from AW in 2013 and 2015, and
similar variability ranges of CDOM properties were mentioned by Pavlov et
al. (2017) north of Svalbard. Therefore, PW in the eastern Fram Strait has
not been advected from the central Arctic Ocean, as in the EGC (Granskog et
al., 2012; Pavlov et al., 2015), but rather it is a modified AW, strongly
affected by heat loss and diluted by sea ice melt in the Barents Sea.
Similar processes also occur on the northern Spitsbergen shelf, where PW was also
found near the ice edge in surface waters diluted and cooled by sea ice
melt.

According to Aas and Høkedal (1996) freshwater runoff from different
sources influence Svalbard waters and there is no universal relation between
salinity and CDOM in this area. Average values of aCDOM(350) in
2014 in PSW (Table 2) were similar to Arctic waters north of the Polar
Front in the Barents Sea described by Hancke et al. (2014) and slightly higher
than observed in this study in 2013 (0.32 ± 0.16 m−1) and 2015
(0.26 ± 0.09 m−1). According to Hancke et al. (2014) the CDOM
pool in the Barents Sea was predominantly of marine origin, while several
studies show terrestrial CDOM in the PW of the EGC (Granskog et al., 2012; Pavlov
et al., 2015), and aCDOM(350) reported for PW in the EGC was
significantly higher, by a factor of 2, than values reported in this study around
Svalbard.

CDOM absorption in WSC reported by Pavlov et al. (2015) and our observations
enabled us to observe significant interannual variability in
aCDOM(350) since 2009 until 2015. The year-to-year changes in
average aCDOM(350) may differ in AW by as much as 200 %
(Table 2). We link these changes with intensity of AW transport to the West
Spitsbergen Shelf presented as spatially and vertically average salinity and
temperature time series (Walczowski et al., 2017). According to this study
the average temperature north of 74∘ N was higher in 2009 than in
2010 and corresponded to lower aCDOM(350) in 2009 relative to
2010 (Pavlov et al., 2015). Similarly in 2013, with the highest CDOM absorption
in our observations, the temperature was lower than in 2014 and 2015
(Walczowski et al., 2017). The average salinity of 35.05 reported in 2014 by
Walczowski et al. (2017) was close to the record high of 35.08 measured in the
period 2000–2016. In 2014 we observed the lowest aCDOM(350)
reported since 2009.

S300–600 varied very little between water masses in a given
season (Table 2); thus we assume that average seasonal values are
representative for all water masses (Table 3). The largest variation in
S300–600 (Fig. 6, Table 3) was observed in 2014, while the
lowest variation in this parameter and a shift towards lower values was
observed in 2013 and 2015. Spectral slope coefficient values
(19.0 ± 2.7 µm−1) reported by Granskog et al. (2012) for
AW across a section in the eastern Fram Strait were very similar to those found
during AREX2013 and AREX2015 (Table 2). Spectral slopes presented by Granskog
et al. (2012), however, were calculated in the broader spectral range
300–650 nm, while Hancke et al. (2014) calculated a spectral slope
coefficient in the narrower spectral range of 350–550 nm. Recalculation of the
spectral slope coefficient for our data set in the spectral range
300–650 nm resulted in an average increase in S by
< 1 µm−1 relative to S300–600. The spectral
slope reported by Hancke et al. (2014) varied among seasons; values in
May 2008 (16 ± 4 µm−1) were higher than those observed
in August 2007 (14 ± 4 µm−1) but both were similar to
values reported in this study. Although Hancke et al. (2014) calculated
spectral slope coefficient for a narrower spectral range, resulting
consistently in lower spectral slope values by ∼ 2 µm−1,
their values were within the range of S300–600 in the current
data set. In the WSC the S300–600 values were higher than those
for surface waters north of Svalbard in winter–spring reported by Kowalczuk
et al. (2017). Observations reported by Kowalczuk et al. (2017) were
conducted earlier in the season and samples were collected below sea ice;
thus
CDOM was less exposed to solar radiation and was potentially less affected by
photobleaching. The highest S300−600 values were found during AREX2014
(20.71 ± 5.26 µm−1), when over 79 % of samples were
classified as AW, which could be associated with photomineralization of DOM in
aging seawater (Obernosterer and Benner, 2004).

4.2 Identification of CDOM sources

According to Stedmon and Markager (2001) the nonlinear relationship between
spectral slope S300–600 and aCDOM(375) allows the
differentiation between terrestrial (allochthonous) and marine (autochthonous)
CDOM pools as well as the assessment of changes in CDOM composition. This approach was
validated by Granskog et al. (2012), who found that CDOM samples taken in PW
with high fractions of meteoric water (i.e., river water) in the western part
of Fram Strait were outside the Stedmon and Markager (2001) model limits for
marine CDOM. Increasing spectral slopes and decreasing CDOM absorption
provides information about degradation of autochthonous CDOM originated from
marine environments (Stedmon and Markager, 2001; Whitehead and Vernet, 2000).
We found decreasing S300–600 values with increasing CDOM
absorption in all three years (Fig. 6). This is similar to that presented by
Kowalczuk et al. (2006) in the Baltic Sea and Pavlov et al. (2014) in
Kongsfjorden, West Spitsbergen. In our study almost all data points are
within the Stedmon and Markager (2001) model limits (Fig. 6) and suggest a
dominant marine (autochthonous) source of CDOM. The highest
S300–600 (> 25 µm−1) with very low CDOM
absorption (< 0.075 m−1) suggests a highly degraded CDOM pool in
2014. In contrast, lower values of S300–600
(< 18 µm−1) with higher absorption (> 0.15 m−1)
could indicate freshly produced CDOM. Lack of correlation between salinity
and aCDOM(λ) was found here (not shown) as by Hancke et
al. (2014), which further suggests a marine origin of organic matter in the
study area.

There were some data points, measured in 2013 characterized by absorption
(> 0.25 m−1) and a spectral slope of ∼ 18 µm−1 that
were outside the upper Stedmon and Markager (2001) model limits. These points
could bias the S300–600 and aCDOM(375) relationship
derived for the present data set, and suggest either a more terrestrial
contribution at high aCDOM(375) from local sources or the influence
of polar water in the western part of the Fram Strait or recirculating
modified AW. A slight increase in humic-like DOM fluorescence
(ICH1 and ICH2), observed near the southwestern
Spitsbergen shelf (Fig. 4), could indicate a small local contribution from a
terrestrial CDOM source.

The presumed molecular structure of marine autochthonous DOM is composed
mainly with low-molecular-weight aliphatic organic compounds characterized
by low saturation with aromatic rings (Harvey et al., 1983). SUVA254
defined by Weishaar et al. (2003) is related to aromatic ring content within
the mixture of water-soluble organic DOM. Massicotte et al. (2017) presented
the global distribution of SUVA254 and found that SUVA254
decreased sharply in the aquatic continuum from fresh
(4.8 m2 gC−1) to oceanic waters (1.68 m2 gC−1).
SUVA254 also decreases with increasing salinity, decreasing rapidly in the salinity
range of 0–8.7, remaining stable at salinity of 8.7–26.8, and decreasing slowly
until salinity reaches oceanic values, and further remaining at a stable level
of ca. 1.7 m2 gC−1 (Massicotte et al., 2017). SUVA254 values
presented in this study (Table 2) were at the lower end of the global range,
close to the oceanic end-member values. The highest average SUVA254
values were found in PSWw in 2013 (1.95 ± 0.60 m2 gC−1) and
PSW in 2014 and 2015 (1.96 ± 0.63 and
1.99 ± 0.30 m2 gC−1, respectively) and the lowest in PSW
(1.31 ± 0.28 m2 gC−1) and AW
(1.41 ± 0.24 m2 gC−1) in 2013 and 2015, respectively.
Pavlov et al. (2016) reported SUVA254 values at a salinity > 34.3 in
the southern Barents Sea waters in the range of 1.3–1.8 m2 gC−1,
which agree well with our findings. The SUVA254 values observed in the
Siberian Shelf at a salinity > 30 varied between
1.25 and 2.3 m2 gC−1 (Gonçalves-Araujo et al., 2015). Low
SUVA254 values suggested overall low saturation of CDOM with aromatic
rings, which supports the hypothesis of predominantly autochthonous CDOM origin
and minor influence by terrestrial DOM in the Nordic Seas, with hydrography
dominated by AW inflow.

4.3 Relationship between CDOM absorption and DOC

The significant amount of DOC in the Arctic Ocean mainly originates from
riverine inflow and permafrost thaw (Stedmon et al., 2011; Amon et al., 2012;
Spencer et al., 2015). The riverine input can be monitored by optical methods
with absorption, fluorescence, or remote-sensing measurements (Spencer et
al., 2012; Walker et al., 2013; Fichot et al., 2013; Mann et al., 2016). The
largest DOC concentrations were found in the Siberian rivers Lena –
1300 µmol L−1, Yenisey – 842 µmol L−1, and
Ob – 950 µmol L−1, and the concentrations were lower in the
North American Yukon – 816 µmol L−1 and McKenzie –
648 µmol L−1 rivers (Amon et al., 2012; Mann et al.,
2016). Both CDOM and DOC in coastal areas in the Arctic Ocean show an inverse
relationship with salinity (Amon et al., 2012) and a very good correlation
between CDOM absorption and DOC has been reported for regions influenced by
riverine input (Matsuoka et al., 2012, 2013; Gonçalves-Araujo et al.,
2015; Pavlov et al., 2016; Mann et al., 2016). The DOC concentration observed
by Amon et al. (2003) in the EGC in the western part of Fram Strait and in
the Denmark Strait was considerably lower and ranged from
76 µmol L−1 in PSW to 55 µmol L−1 in AW. Amon
et al. (2003) found a weak inverse relationship between DOC and salinity in
the Nordic Seas and a weak correlation between DOC and CDOM fluorescence. The
DOC concentration reported in this study in the AW-dominated eastern part of
Fram Strait was similar to that reported by Amon et al. (2003) in the EGC but
lower than found in Barents Sea waters entering the White Sea at salinities
close to 34.9 (Pavlov et al., 2016). The DOC concentration in the open Laptev
Sea was over 100 µmol L−1 as reported by Gonçalves-Araujo
et al. (2015). We observed a very weak correlation between DOC concentration
and aCDOM(350) (Fig. 7). That could be explained by the low
number of samples influenced by terrestrial humic substances in our data,
which have elevated aCDOM(350), DOC, and lower salinity.
Additionally, our data were at the lower range of the globally observed
distribution of DOC and aCDOM(350), where the relationship is
characterized by large uncertainty (Massicotte et al., 2017).

Figure 7Relationship between aCDOM(350) and DOC and linear
relationship between those parameters in 2013–2015.

The relationship between the carbon-specific CDOM absorption coefficient
aCDOM*(350) and S275–295 was another approach
suggested by Fichot and Benner (2011, 2012) in the Gulf of Mexico to trace
the influence of terrigenous DOC in coastal margins and to estimate DOC from
optical measurements. We did not observe a significant relationship between
aCDOM*(350) and S275–295 (not shown). However,
aCDOM*(350) as a function of S300–600 showed much
more promise (Fig. 8). This could be potentially applied for DOC estimations
from CDOM absorption measurements in Nordic Seas.

4.4 Distribution of FDOM components in the ocean and their dependence on allochthonous and autochthonous sources

The distribution pattern of the main FDOM components with depth in the global
oceans' biogeochemical provinces is significantly different for humic-like
and protein-like FDOM (Stedmon and Nelson, 2015; Catalá et al., 2016).
The intensity of the humic-like FDOM fraction is usually higher close to
continental margins and significantly depleted in the centers of subtropical
gyres (Murphy et al., 2008; Jørgensen et al., 2011; Kowalczuk et al.,
2013). The fluorescence of humic-like DOM fractions is low in the surface
layer and rapidly increases with depth, reaching a constant high level
below 200 m. Protein-like FDOM fluorescence intensity usually
increases toward the open ocean and the highest intensity is observed in the
surface waters, rapidly decreasing with depth, reaching a constant low level
below the epipelagic layer (Jørgensen et al., 2011; Kowalczuk et al., 2013;
Catalá et al., 2016). Such profiles indicate that amino-acid-like DOM is
linked to surface water production. Catalá et al. (2016) demonstrated
that the global depth distribution tryptophan-like FDOM component has a
local maximum associated with a chlorophyll a maximum. The linkage between
protein-like components and chlorophyll a concentration shown
qualitatively in the global ocean by Stedmon and Nelson (2015) and Catalá
et al. (2016) was previously confirmed quantitatively in mesocosm studies, for example, Romera-Castillo et al. (2010), which indicated that
phytoplankton excreted tryptophan-like fluorophores, and tryptophan-like
component concentration has been related to primary production (Brym et al.,
2014). In situ quantitative correlation between chlorophyll a
concentrations and fluorescence intensity of the protein-like FDOM fraction has
been observed and documented recently. Yamashita et al. (2017) reported
significant positive correlation between the tryptophan-like component and
Chla
(r= 0.53, p< 0.001) in the surface waters of the Pacific Ocean.
Yamashita et al. (2017) also found spatial coupling between the
tryptophan-like component and chlorophyll a concentration, which was
strongest in the Bering Sea. A study by Loginova et al. (2016) from a Peruvian
upwelling system also reported a positively correlated chlorophyll a
concentration and protein-like component (R2= 0.40,
p< 0.001).

The distribution of fluorescence intensity of the main FDOM components in the
Nordic Seas, dominated by warm water of Atlantic origin, followed the general
trends observed globally. The highest FDOM intensity, especially of
humic-like components, was observed close to continental margins, in the
vicinity of major river outflows. Para et al. (2013) observed significant
inverse trends of humic-like FDOM components with salinity in the Canadian
shelf of the Beaufort Sea close to McKenzie River outflow. Similar
observations were documented by Gonçalves-Araujo et al. (2015) in the
Lena River delta at the Siberian Shelf and by Pavlov et al. (2016) near the
Northern Dvina River outlet in the White Sea. The impact of the humic-like
FDOM component on DOM composition decreased with increased distance from
freshwater sources and increased salinity, where the protein-like FDOM
fraction became dominant, for example, outside of the McKenzie River plume in
the Beaufort Sea (Para et al., 2013) and in the White Sea (Pavlov et al.,
2016). In the Fram Strait the distribution of humic-like fluorescence
(ex. = 340 nm and em. = 420 nm) observed by Amon et
al. (2003) in the Fram and Denmark straits was related to large-scale water
mass distribution in the Nordic Seas and was characterized by elevated values
of FDOM intensity in the western part of Fram Strait that was under the
influence of the EGC, and characterized by low FDOM intensity and FDOM
intensity uniformly distributed with depth in the core of AW in its eastern
part. The FDOM distribution in AW shown by Amon et al. (2003) corresponded
well to vertical profiles of ICH1 and ICH2 in AW,
shown in Fig. 4. This was also in good agreement with CDOM distribution in
the Fram Strait (Granskog et al., 2012; Pavlov et al., 2015) and FDOM
humic-like fraction (ex. = 280 nm and em. = 450 nm) distribution presented by
Granskog et al. (2015b). The humic-like fraction of DOM in the eastern Fram
Strait is more than 10 times lower compared to PW in the EGC (Granskog et
al., 2015b). A 20 m layer of less saline water diluted by sea ice melt
characterized by significantly lower humic-like FDOM intensity overlaid the
PW water with high FDOM intensity in the EGC (Granskog et al., 2015b).

In situ fluorometry provided an opportunity to study FDOM distribution in
greater detail and commercially available FDOM fluorometers are usually built
to detect humic substances (Amon et al., 2003; Belzile et al., 2006;
Kowalczuk et al., 2010; Aiken et al., 2011; Loginova et al., 2016). In this
study we simultaneously measured three different FDOM components, and the
most interesting feature observed with the use of this new instrument was very
significant spatial coupling between ICH3 and
IFChla. Similarities in the vertical distribution of
protein-like FDOM, ICH3, stimulated chlorophyll a fluorescence intensity, IFChla, and total non-water
absorption coefficient at 676 nm, atot−w(676), implied
quantitative interrelation among those parameters and the same dominant factor
controlling these parameters in time and space. We found a significant
positive correlation (R2= 0.65, p< 0.0001) between
ICH3 and IFChla (Fig. 5a), which suggests that
production of protein-like FDOM is closely related to spatial and temporal
phytoplankton dynamics. Additionally, a statistically significant dependence
of ICH3 and Chla concentration from water samples indicated
that phytoplankton biomass is an important source of protein-like FDOM.

We observed significant interannual variation in CDOM optical properties in
the Nordic Seas. It is likely that these year-to-year changes in CDOM
absorption coefficient and spectral slope coefficient were related to the
intensity of AW inflow to the Nordic Seas. According to Walczowski et al. (2017)
there was very strong interannual variability in AW inflow overlaid on the
long-term increasing trend. CDOM absorption decreased and spectral slope
coefficient increased during years when an increase in temperature was observed
for AW (Walczowski et al., 2017), e.g., in 2009 (Pavlov et
al., 2015) and in 2014 (this study). Decrease in AW temperature was
accompanied by mutual increase in aCDOM(350) and decrease in
S300–600, e.g., in 2010 (Pavlov et al., 2015) and in 2013 and
2015 (this study). We surmise that during less intense inflow of AW to the Nordic
Seas a higher proportion of PW is transported with the ESC and SC to the eastern part
of Fram Strait, contributing to the increase in CDOM in West Spitsbergen Shelf
waters.

In situ observations with the use of a three-channel fluorometer coupled with
other optical instruments enabled us to show a significant correlation
between protein-like FDOM and chlorophyll a in the Nordic Seas.
Quantitative dependence between protein-like FDOM (ICH3) and
chlorophyll a fluorescence (IFChla) and between
protein-like FDOM (ICH3) and total non-water absorption
coefficient at 676 nm (atot−w(676)) based on direct in situ
observations clearly indicated that phytoplankton biomass is the primary
source of low-molecular-weight DOM fraction in Nordic Seas influenced by warm
AW. This highlights the role of phytoplankton dynamics as an important factor
controlling FDOM/CDOM. The
freshly produced protein-like FDOM fraction did not contribute to CDOM/FDOM
optical properties observed in the visible spectral range as its fluorescence
excitation (absorption) and emission characteristics were located in the
ultraviolet spectral range. Observed variability in spectral index
(aCDOM*(350), SUVA254, S300–600) values
suggest that CDOM/FDOM in the Nordic Seas has an autochthonous origin. Yet,
further investigation of the DOM transformation processes from labile freshly
produced protein-like DOM fractions to more complex organic molecules is
needed to better understand the CDOM/FDOM dynamics in the Nordic Seas.
Typically humic-like FDOM was found in low concentrations in the study area,
showcasing the limited terrestrial influence, in contrast to the East
Greenland Current, for example (Gonçalves-Araujo et al., 2016).

Dissolved organic carbon (DOC) was weakly correlated with
aCDOM(350) in the study area, likely due to limited terrestrial
influence, and aCDOM(350) shows no promise to be used as a tool
to predict DOC. The same was the case for spectral slope at short wavelengths
(S275–295), proven earlier to work for nearshore environs
(Fichot and Benner, 2011, 2012). Conversely, there was a significant
inverse nonlinear relationship of CDOM-specific DOC absorption
(aCDOM*(350)) with spectral slope at a broader spectral range
(S300–600). This relationship provides a potential for indirect
estimates of DOC with the use of optical measurements in this region.

All data used in this study will be freely available, for
scientific use only, upon request. Anyone interested in using this data set
for scientific research should contact the corresponding author via e-mail.

We thank the crew of R/V Oceania and colleagues for the help
onboard. This work was supported by the Polish–Norwegian Research Programme
operated by the National Centre for Research and Development under the
Norwegian Financial Mechanism 2009–2014 in the frame of project contract
Pol–Nor/197511/40/2013, CDOM–HEAT. This work was partially financed from
the funds of the Leading National Research Centre (KNOW) received by the
Centre for Polar Studies for the period 2014–2018. Mats A. Granskog was
supported by the Centre for Ice, Climate and Ecosystems (ICE) at the
Norwegian Polar Institute, and Alexey K. Pavlov by the Research Council of
Norway through the STASIS project
(221961/F20).